impact of millennial-scale holocene climate variability on...

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Global and Planetary Chan

Impact of millennial-scale Holocene climate variability on

eastern North American terrestrial ecosystems:

pollen-based climatic reconstruction

Debra A. Willarda,*, Christopher E. Bernhardta,

David A. Korejwoa, Stephen R. Meyersb

aU.S. Geological Survey, 926A National Center, 12201 Sunrise Valley Drive, Reston, VA 20192, United StatesbGeology and Geophysics Department, Yale University, P.O. Box 208109, New Haven, CT, 06520-8109, United States

Received 17 March 2004; accepted 30 November 2004

Abstract

We present paleoclimatic evidence for a series of Holocene millennial-scale cool intervals in eastern North America that

occurred every ~1400 years and lasted ~300–500 years, based on pollen data from Chesapeake Bay in the mid-Atlantic region

of the United States. The cool events are indicated by significant decreases in pine pollen, which we interpret as representing

decreases in January temperatures of between 0.28 and 2 8C. These temperature decreases include excursions during the Little

Ice Age (~1300–1600 AD) and the 8 ka cold event. The timing of the pine minima is correlated with a series of quasi-periodic

cold intervals documented by various proxies in Greenland, North Atlantic, and Alaskan cores and with solar minima

interpreted from cosmogenic isotope records. These events may represent changes in circumpolar vortex size and configuration

in response to intervals of decreased solar activity, which altered jet stream patterns to enhance meridional circulation over

eastern North America.

D 2004 Elsevier B.V. All rights reserved.

Keywords: Paleoclimatology; Holocene; Little Ice Age; Chesapeake Bay; pollen

1. Introduction

Debates on the causes of millennial-scale climate

variability during an interglacial have included solar

forcing, internally forced changes in deep oceanic

0921-8181/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.gloplacha.2004.11.017

* Corresponding author. Tel.: +1 703 648 5320; fax: +1 703 648

6953.

E-mail address: [email protected] (D.A. Willard).

circulation, and modulations of atmospheric circula-

tion such as the North Atlantic Oscillation (Bond et

al., 2001; Keigwin and Pickart, 1999; Mayewski et

al., 1997; Shindell et al., 2001). Such variability has

been documented in Holocene records of drift ice in

the North Atlantic Ocean (Bond et al., 2001), d18O in

Irish speleothems and Greenland ice cores (Johnsen et

al., 2001; McDermott et al., 2001), sea-surface

temperatures in the subtropical southeastern Atlantic

ge 47 (2005) 17–35

http://www.ngdc.noaa.gov/paleo/pollen.html

D.A. Willard et al. / Global and Planetary Change 47 (2005) 17–3518

Ocean (deMenocal et al., 2000), and biogenic silica

and pollen records from Alaskan lakes (Hu et al.,

2003). Given recent concerns regarding the impact

of anthropogenic greenhouse gases on global cli-

mate, this broad-scale signal of quasi-periodic

(1500F500 years) climate forcing needs to be

understood in the context of its influence on natural

trends in global temperatures, precipitation, and sea-

ice extent over the next few centuries. Because

climate response to external forcing factors may not

be globally uniform or synchronous and because

feedbacks exist among different forcing factors, it is

important to develop an understanding of their

relative roles in driving climate periodicity. To do

so, examination of high-resolution paleoclimate

records from different regions and ecosystems is

critical. Thus far, Holocene records clearly docu-

menting millennial-scale climate variability have

been confined primarily to marine, ice core, and

high-latitude lacustrine sites; no Holocene records

have been available from mid-latitude terrestrial

sites in the eastern United States with sufficient

temporal resolution to examine millennial-scale

varibiality.

In this paper, we present high-resolution sediment

records from Chesapeake Bay, which detail the impact

of Holocene climate variability on terrestrial and

estuarine ecosystems in a key climatic region of

North America. Chesapeake Bay is the largest estuary

in the United States, extending N300 km in length,

covering an area of 6500 km2, and draining a

watershed of 166,000 km2 (Fig. 1). The bay formed

when rapid sea-level rise between ~8.0 and 7.6 ka

flooded the paleo-Susquehanna River system (Cronin

et al., 2003; Hobbs, 2004), abruptly changing the

fluvial system to an estuary. Chesapeake Bay has a

deep (20–50 m) channel formed by the Susquehanna

River during sea-level lowstands of the last glacial

period and a thick (10 to N20 m) Holocene sequence

identified by geophysical surveys (Colman and Halka,

1989; Halka et al., 2000; Vogt et al., 2000). Pre-

Colonial Holocene sedimentation rates of 0.2–1.0 cm

year�1 provide a stratigraphic record of paleoclimatic

variability with temporal resolution that exceeds most

lacustrine and marine records. This record includes an

almost complete Holocene record and abundant floral

(Willard et al., 2003), faunal (Cronin et al., 2000,

2003), and geochemical (Zimmerman and Canuel,

2000) proxies to evaluate ecosystem response to

climatic fluctuations on a variety of temporal scales.

Chesapeake Bay lies within the Atlantic Coastal

Plain and is surrounded by oak–pine forests in the

southern two-thirds of the bay (south of ~388Nlatitude); oak–chestnut and tulip poplar forests dom-

inate along the northern reaches of the bay (Braun,

1950; Brush et al., 1980; Greller, 1988). The boundary

between the two forest types corresponds to changes in

substrate type, drainage potential of the native soils,

and climatic regime. Oak–pine forests occupy uncon-

solidated sediments with low topographic relief in the

coastal plain; deciduous forests occupy drier sites with

greater topographic relief to the north. South of the

Chesapeake Bay mouth, pines become progressively

more abundant and grade into the southeastern pine

forests (Braun, 1950; Greller, 1988).

A climatic transition zone delineated by the

position of the polar jet stream crosses Chesapeake

Bay, and changes in its position directly influence

regional temperature, precipitation, and storm fre-

quency and intensity (Vega et al., 1998, 1999). The

position of the jet stream generally reflects the shape

and size of the circumpolar vortex, which incorporates

atmospheric circulation patterns in both eastern and

western hemispheres. During times of meridional

circulation, the jet stream exhibits an enhanced ridge

and trough configuration, whereas a flattened jet

stream characterizes intervals of zonal circulation. In

the eastern United States, intervals of enhanced

meridional flow, such as occurred between 1966 and

1990, result in the advection of polar air south, cooler

temperatures, and increased cyclone frequency (Davis

and Benkovic, 1992). The impact of relatively subtle

shifts in jet stream position on temperature and

precipitation in the mid-Atlantic region makes Ches-

apeake Bay an ideal site for paleoclimate studies.

The possible influence of solar variability on

atmospheric circulation patterns and the earth’s climate

recently has been explored through analysis of satellite

monitoring data, which have a limited period of record

from the mid-1940s to the present (Angell, 1992;

Davis and Benkovic, 1992), and using historical

compilations of sun spot records, which extend

observational records to 2000 years BP (Pang and

Yau, 2002). These studies show correlations between

solar cycles, circumpolar vortex configuration, and

various atmospheric circulation patterns (i.e., El Nino,

BALTIMORE

WASHINGTON

RICHMOND

2209

2208

2207

DELAWARE BAY

0 km 40

PO TOMACRIV

ER

77°W 76°W 75°W

39°N

38°N

37°N

PTMC 3,

ATLA

NTI

C O

CEA

N

CH

ES

AP

EA

KE

BAY

Meridional

MeanZonal

CB

A

B

C

D

E

F

GI

H

Fig. 1. Location of coring sites in Chesapeake Bay (CB) in eastern North America. Core sites include MD99-2207, MD99-2208, MD99-2209,

and PTMC3-2. Lettered circles indicate collection sites for samples used to test validity of 400-year marine correction for 14C dates. Circles

A–C represent location of modern oyster samples (Colman et al., 2002); circles D–I represent location of cores in which shells from post-

Colonial samples were radiocarbon dated in this study (see text for explanation). Inset map shows position of Chesapeake Bay on eastern coast

of North America, a simplified illustration of the mean position of the polar jet stream (thick line), and its altered configuration during zonal

(dashed line) and meridional (thin line) flow regimes.

D.A. Willard et al. / Global and Planetary Change 47 (2005) 17–35 19

quasi-biennial oscillation) over decadal to centennial

time scales (Angell, 1998; Shindell et al., 2001).

Although only a few mechanisms for solar control on

atmospheric circulation patterns have been proposed

(i.e., Tinsley, 2000), the apparent correlations support

earlier hypotheses that the size and intensity of the

vortex, perhaps influenced by solar variability, may

have governed much post-Glacial climate variability

on centennial to millennial time scales (Willett, 1949).

The decadal- to centennial-scale resolution of the

Chesapeake Bay sedimentary record provides an

opportunity to evaluate the impact of such variability

on a terrestrial mid-latitude record. Pollen is well

preserved throughout the sediments, and calibration of

pollen assemblages from surface sediments with plant

assemblages in adjacent forests along the east coast

facilitates interpretation of changes in forest composi-

tion, temperature, and precipitation.

2. Calibration of pollen and climatic data

Pollen assemblages preserved in sediments from

North American estuaries and the western Atlantic

Pin

us p

erce

nt

Annual Precipitation100-130 cm

Annual Precipitation130-180 cm

B)

25

30

35

40

450 20 40 60 80 100

Latit

ude

(˚N

)

Pinus percent

A) x = -5.422y + 250.356 r 2 = 0.887

x = 0.221y - 7.225 r 2 = 0.910

20151050-5

0

20

40

60

80

100

January temperature (˚C)

Fig. 2. (A) Regression model (r2=0.887) for Pinus pollen

abundance in estuarine and shallow marine sediments of the

western Atlantic Ocean vs. latitude. (B) Regression mode

(r2=0.910) for Pinus pollen abundance in estuarine and shallow

marine sediments of the western Atlantic Ocean vs. January

temperature. Pollen data compiled from Buening et al. (2000)

Edwards and Willard (2001), Litwin and Andrle (1992), and Willard

(new data available on North American Pollen Database: http:/

www.ngdc.noaa.gov/paleo/pollen.html). Climate data were taken

from nearest weather stations and averaged over the years 1971–

2000 (National Climate Data Center, 2000).


Ocean shelf provide a regional record of vegetation

on the adjacent coasts because the prevailing westerly

breezes during the flowering season deposit pollen

abundantly in shallow marine and estuarine sediments

(Mudie, 1980, 1982; Litwin and Andrle, 1992;

Willard, 1994). To assess the correlation between

abundance of pollen from dominant wind-pollinated

taxa in the region and climatic parameters, we

compared Pinus and Quercus pollen abundance in

estuarine and shallow marine (V100 m) sediments

from the western Atlantic Ocean (Emery, 1966) with

latitude, mean January temperature, and mean annual

precipitation from the nearest climate monitoring

stations. Climatic data were averaged over the years

1971–2000 (NCDC, 2000). Pollen data were obtained

from Buening et al. (2000), Litwin and Andrle

(1992), Edwards and Willard (2001) and Willard

(new data available online at the North American

Pollen Database (NAPD) at the World Data Center

for Paleoclimatology in Boulder, CO: http://

www.ngdc.noaa.gov/paleo/pollen.html).

Abundance of Pinus pollen in modern sediments

shows a clear south to north gradient of high to low

percent abundance (Fig. 2A), ranging from N80% at

288N to b20% at 428N. This gradient mirrors the

abundance of southern pines (Pinus echinata, Pinus

elliottii, Pinus palustris, Pinus taeda, Pinus virgin-

iana) in eastern United States forests. Their approx-

imate northern limit is ~388N (Bartlein et al., 1986);

measurements of current timber stock indicate that

southeastern states (Virginia to Florida) contain a

combined stock of 45,147 million cubic feet (mcf) of

southern pines, compared to 1493 mcf from Maryland

through the New England states (Smith et al., 1997).

These pine species require warm winters (N�0.6 8C),high moisture availability (N885 mm annual precip-

itation), and unconsolidated soils (Iverson et al., 1999;

Thompson et al., 2000). Quercus pollen abundance

shows no correlation with latitude or climatic param-

eters (r2b0.011), which is consistent with their

common distribution throughout eastern North Amer-

ican forests.

Pinus pollen abundance and mean January temper-

atures are strongly correlated (r2=0.91; Fig. 2B),

providing a proxy for winter temperature fluctuations.

The much weaker correlation between Pinus pollen

abundance and mean annual precipitation (r2=0.331)

reflects the lack of a strong latitudinal precipitation

l

,

/




gradient. However, samples collected north of 378latitude were comprised of b50% Pinus pollen, and

those sites receive an average of 100–130 cm annual

precipitation, whereas those collected south of 368Nare characterized by N65% Pinus pollen and receive

130–180 cm annual precipitation (Fig. 2A). This

relationship is consistent with data on Pinus pollen

production rates in the Washington, DC region, which

covary with precipitation of the previous year. Over

an 11-year period, comparison of atmospheric pollen

concentrations with June to December precipitation

(when cone production occurs) indicates that precip-

itation during that period explains more than one-third

of the variation in atmospheric Pinus pollen abun-

dance. When precipitation is greatest, atmospheric

Pinus pollen concentrations are greatest (Willard et

al., 2003). Therefore, we use Pinus pollen abundance

as a proxy for both winter temperature and mean

annual precipitation. Using only the relationship

between pollen abundance and winter temperature

derived above, a 20% increase in pine pollen in

estuarine and shallow-marine sediments is equivalent

to an increase in January temperature of ~4.5 8C (Fig.

2B). However, because of the combined influences of

winter temperature and precipitation, it is likely that

temperature reconstructions based only on the pine–

temperature correlation overestimate absolute temper-

ature variability to some extent.

3. Holocene pollen record from Chesapeake Bay

sediment cores

3.1. Stratigraphy and age

The Holocene vegetational history presented herein

is based on pollen assemblages from four sediment

cores collected in the mainstem of Chesapeake Bay

(Fig. 1). We used a modified Calypso corer on the

IMAGES V cruise of the R/V Marion Dufresne to

collect three cores in 1999: MD99-2207 (38801.83VN,76812.88VW, 2070 cm long, 25 m water depth

[mwd]); MD99-2208 (38832.24VN, 76829.19VW, 782

cm long, 10 mwd); and MD99-2209 (38853.18VN,76823.68VW, 1720 cm long, 26 mwd). We collected a

shorter piston core (PTMC3-2, 38801.6118VN,

76813.1938VW, 414 cm long, 23.1 mwd) on the R/V

Kerhin in 1996.

Forty-three accelerator mass spectrometer (AMS)

radiocarbon dates (all dates are presented as calibrated

years before present) (Table 1) and well-documented

biostratigraphic events from numerous eastern North

American sites form the basis for age models for these

cores. Age models for each core were determined by a

combination of linear interpolation and best-fit poly-

nomials between radiocarbon dates obtained on

molluskan shells and foraminifera, calibrated with

the marine reservoir correction of 400 years (Fig. 3).

Use of the marine correction is based on analyses of

oysters collected live in Chesapeake Bay before

atmospheric nuclear testing and shells from sediment

cores collected within the zone identified as 1880–

1910 AD based on pollen biostratigraphy. The oyster

shells yielded an average reservoir correction of

365F143 years (Colman et al., 2002). We also

identified intervals in six sediment cores collected

south of the estuarine turbidity maximum that repre-

sent the interval of peak land clearance in the water-

shed (1880–1910 AD) using the abundance of ragweed

(Ambrosia) pollen (after Willard et al., 2003). Shells of

Mulinia, an unidentified mollusk, and the bryozoan

Haminoea solitaria were collected from the interval

and dated using AMS (Table 2). The mean reservoir

effect from these 10 samples is 374.5F60 years, also

comparable to the marine correction. These data

indicate that the standard marine reservoir correction

of 400 years is appropriate for use in Chesapeake Bay

sites south of the estuarine turbidity maximum.

The basal date of 10.5 ka in core MD99-2207 is

based on percent abundance of pine (Pinus) and oak

(Quercus) (23% and 53%, respectively). Such high

abundances in eastern North American pollen assemb-

lages were not attained until at least 10.5 cal ka

(Kneller and Peteet, 1993; Watts, 1979); therefore,

this is a conservative estimate, and basal sediments

may be younger. Confirmation of this estimate is

provided by comparison of paleomagnetic inclination

data from MD99-2207 with the composite inclination

curve from the northeastern United States, which

places the base of the core at 10–10.5 ka (Clifford

Heil, personal communication, 2003). The Carya

increase abundance from b2% to N4% at 1300–1250

cm in MD99-2207 correlates with an increase at ~9.4

ka documented in mid-Atlantic ponds and bogs

(Watts, 1979) and is within the error range of the

radiocarbon date of 9.21 (9.63–9.04) ka at 1250 cm.

Table 1

Locality, water depth, and age information for cores MD99-2209, PRCK 1-2, PTXT 2-P-5, and PTMC 3-P-2

Site and

core no.

Latitude

(8N)Longitude

(8W)

Water

Depth (m)

Lab # Depth of 14C

date (cm)

Material

dated

d13C 14C age

(conventional)

1ra Calibrated age

(years BP)

F2rb

(years BP)

Rhode River

MD99-2209 38853.18V 76823.68V 26 OS-21226 296 bivalve �0.87 610 30 270 300–150

RD 98-1c 38853.202V 76823.502V 26.5 OS-19215 340 bivalve �0.87 725 55 340 270–460

MD99-2209 38853.18V 76823.68V 26 OS-21381 369 bivalve �0.57 745 35 410 450–510

MD99-2209 38853.18V 76823.68V 26 OS-21382 455 bivalve �0.68 1150 40 680 760–640

RD 98-1c 38853.202V 76823.502V 26.5 OS-19214 457 bivalve �1.03 1150 85 675 550–870

MD99-2209 38853.18V 76823.68V 26 OS-21227 485 bivalve �1.29 1240 30 770 870–710

MD99-2209 38853.18V 76823.68V 26 OS-21383 576 bivalve �0.9 1600 35 1165 1230–1070

MD99-2209 38853.18V 76823.68V 26 OS-21384 665 bivalve �1.73 2050 40 1610 1700–1520

MD99-2209 38853.18V 76823.68V 26 OS-21228 733 bivalve �0.77 2210 35 1810 1880–1720

MD99-2209 38853.18V 76823.68V 26 OS-21229 780 bivalve �0.74 2500 35 2140 2280–2070

MD99-2209 38853.18V 76823.68V 26 OS-21385 820 bivalve �0.18 4230 40 4340 4210–4420

MD99-2209 38853.18V 76823.68V 26 OS-21230 904 bivalve �0.7 5530 40 5905 5810–5990

MD99-2209 38853.18V 76823.68V 26 OS-21231 1029.5 bivalve �0.14 5690 40 6100 5980–6180

MD99-2209 38853.18V 76823.68V 26 OS-21232 1159 bivalve �0.08 5690 40 6380 6290–6440

MD99-2209 38853.18V 76823.68V 26 OS-21233 1199 bivalve 0.02 5980 40 6390 6300–6460

MD99-2209 38853.18V 76823.68V 26 OS-21488 1439 bivalve �0.74 6250 35 6700 6620–6770

MD99-2209 38853.18V 76823.68V 26 OS-21386 1605 bivalve �3.73 6290 35 6730 6650–6830

MD99-2209 38853.18V 76823.68V 26 OS-21489 1694 bivalve �3.53 8670 45 9220 9050–9660

MD99-2209 38853.18V 76823.68V 26 OS-21387 1694 oyster �1.04 6660 45 7200 7080–7280

MD99-2209 38853.18V 76823.68V 26 OS-21388 1705 bivalve �1.49 7050 40 7550 7450–7593

MD99-2209 38853.18V 76823.68V 26 OS-21389 1720 bivalve �1.5 7100 45 7570 7500–7650

Parker Creek

MD99-2208 38832.2400V 76829.1900V 10 WW 3392 59 mollusk 0 670 40 297 258–414

MD99-2208 38832.2400V 76829.1900V 10 WW-3393 559 bivalve �3 5080 40 5447 5317–5536

MD99-2208 38832.2400V 76829.1900V 10 WW-2705 629 bivalve 0 6055 65 6462 6307–6632

MD99-2208 38832.2400V 76829.1900V 10 WW-3395 751.5 oyster �4.6 6980 40 7410 7316–7465

MD99-2208 38832.2400V 76829.1900V 10 WW-3396 765 oyster �5.7 7360 40 7810 7721–7915

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22

MD99-2208 38832.2400V 76829.1900V 10 WW-3397 779 oyster �8.7 7780 40 8217 8151–8336

MD99-2208 38832.2400V 76829.1900V 10 WW-2706 779 oyster �5.6 7740 80 8177 8004–8360

Potomac River

PTMC 3-P-2 38801.6118V 76813.1938V 23.1 OS 15679 81 Mulinia lateralis 0.01 540 30 150 0–260

PTMC 3-P-2 38801.6118V 76813.1938V 23.1 WW 1284 141 Mulinia lateralis 0.1 540 50 150 0–280

PTMC 3-P-2 38801.6118V 76813.1938V 23.1 OS 15680 161 Mulinia lateralis �0.29 885 35 500 450–540


PTMC 3-P-2 38801.6118V 76813.1938V 23.1 WW 1589 225 Mulinia lateralis 0 990 40 550 510–640

PTMC 3-P-2 38801.6118V 76813.1938V 23.1 IS 17242 229 Elphidium �1.72 1230 30 750 680–830


PTMC 3-P-2 38801.6118V 76813.1938V 23.1 OS 17508 331 Elphidium �2.41* 2450 256 2080 1530–2710

MD99-2207 38801.8300V 76812.880V 25 OS-21487 221.5 bivalve �0.42 855 25 490 450–510

MD99-2207 38801.8300V 76812.880V 25 OS-25825 377.5 bivalve �2.04* 125 50 0 –

MD99-2207 38801.8300V 76812.880V 25 OS-21670 387.5 bivalve 0.11 4100 45 4140 4000–4270

MD99-2207 38801.8300V 76812.880V 25 OS-21671 573.5 bivalve �0.13 4470 45 4630 4530–4790

MD99-2207 38801.8300V 76812.880V 25 OS-25826 687.5 bivalve 0.14 4590 55 4810 4630–4930

MD99-2207 38801.8300V 76812.880V 25 OS-21664 777 bivalve 0.2 6130 55 6560 6420–6670

MD99-2207 38801.8300V 76812.880V 25 OS-21665 833.5 bivalve 0.18 6430 65 6900 6740–7080

MD99-2207 38801.8300V 76812.880V 25 OS-21666 901 bivalve �8.09* 9150 65 9810 9530–10,090

MD99-2207 38801.8300V 76812.880V 25 OS-25827 901 bivalve 0.7 6540 45 7025 6930–7160

MD99-2207 38801.8300V 76812.880V 25 OS-25828 960 bivalve �1.94* 8150 55 8600 8470–8850

MD99-2207 38801.8300V 76812.880V 25 OS-21667 993 bivalve 0.37 7080 60 7560 7450–7650

MD99-2207 38801.8300V 76812.880V 25 OS-25829 1152.5 bivalve �7.27* 8930 65 9475 9100–9820

MD99-2207 38801.8300V 76812.880V 25 OS-21668 1161 bivalve �9.66* 9400 100 10,130 9710–10,560

MD99-2207 38801.8300V 76812.880V 25 OS-21503 2051 wood �27.78* 10,400 45 12,340 11,950–12,800

MD99-2207 38801.8300V 76812.880V 25 B-160832 1045 non-marine mollusc �9.5* 9440 60 7894 7434–7964

MD99-2207 38801.8300V 76812.880V 25 B-106833 1090 non-marine mollusc �9* 9220 40 8191 8041–8311

MD99-2207 38801.8300V 76812.880V 25 B-106834 1249 bivalve �1.6 8800 40 9210 9040–9630

Radiocarbon dates compiled from Cronin et al. (2000, 2003) and Colman et al. (2002). Dates marked with an asterisk were not used in age models because d13C values suggested

transport or reworking from adjacent units or because they represent reversals.a One-sigma counting errors.b Upper and lower limits of range based on two-sigma errors in calibration.c From Zimmerman and Canuel (2000).

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0

1000

2000

3000

4000

5000

6000

7000

8000

9000

0

100

200

300

400

500

600

700

800

MD99-2208

0

1000

2000

3000

4000

5000

6000

7000

8000

0

200

400

600

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1000

1200

1400

1600

1800

MD99-2209

0

1000

2000

3000

4000

5000

6000

7000

8000

9000

1000

0

1100

0

0

200

400

600

800

1000

1200

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2000

MD99-2207D

epth

(cm

)

Calibrated yrBP

0

500

1000

1500

2000

2500

3000

0

100

200

300

400

PTMC 3

Fig. 3. Plot of calibrated radiocarbon ages against depth in cores MD00-2207, MD99-2208, and MD99-2209. Where error bars are not visible,

they are smaller than the boundaries circumscribed by the size of the symbols. Diamonds represent radiocarbon dates used to develop age

models; squares represent dates not used in age models because of age reversals, or d13C values that indicate transport or reworking from

adjacent units. Open diamonds represent dates in the uppermost parts of the cores based on lead-210 analysis and Ambrosia pollen

biostratigraphy. Age models for pre-Colonial (pre-1700 AD) segments of each core are as follows. MD99-2208: 59–559 cm:

age=(10.398*depth)�316.482; 559–629 cm: age=(14.5008*depth)�2658.500; 629–752 cm: age=(7.739*depth)+1594.310. MD99-2207:

387–573 cm: age=(2.634*depth)+3119.167; 573–687 cm: age=(1.579*depth)+3724.474; 687–777 cm: age=(19.553*depth)� 8632.737; 777–

833 cm: age=(6.018*depth)+1884.248; 833–901 cm: age=(1.852*depth)+5356.481; 901–993 cm: age=(5.815*depth)+1785.489; 993–1249

cm: age=(6.334*depth)+1159.805; 1249–2050 cm: age=(1.610*depth)+7198.502. MD99-2209: age model for upper 800 cm provided in

Willard et al., 2003; 819–904 cm: age=(18.631*depth)�10937.381; 900–1200 cm: age=(1.743*depth)+4324.51; 1200–1439 cm:

age=1.292*depth)+4841.292; 1439–1694 cm: age=(1.961*depth)+3878.431. PTMC 3 (160–450 cm): age=(0.012*depth2)+

(0.97*depth)�41.675 (r2=0.86).


The Tsuga decline at 730 cm in MD99-2207 and 840

cm in MD99-2209 is dated at 5.4 ka (Fuller, 1998).

The rise in Ambrosia abundance at 240 cm (MD99-

2209) and 160 cm (MD99-2207) represents peak land

clearance at ~1880–1910 AD (Brush, 1984; Willard et

al., 2003). The greatest uncertainty for these age

models lies within the intervals between 2.2 ka and

4.0 ka and 5.0 ka and 5.8 ka (Fig. 3), which include

Table 2

Comparison of expected and actual radiocarbon dates on shells collected in sediments deposited between 1880 AD and 1910 AD, based on Ambrosia pollen biostratigraphy

Laboratory

number

Core

identification

Sample location Sample

depth (cm)

Material d13C (x)a Age

(14C years BP)

Errorb Estimated

pollen agecExpected 14C

agedReservoir

effecte

Beta 161216f PC6B-2 Pocomoke Sound 202–204 Haminoea

solitaria

�0.9 490 40 1880–1910 AD 86.5F10.8 403.5

Beta 161221f PC2B-3 Pocomoke Sound 268 Mulinia �0.8 460 40 1880–1910 AD 86.5F10.8 373.5

Beta 161222f PC2B-3 Pocomoke Sound 380–386 Mulinia �0.7 440 40 1880–1910 AD 86.5F10.8 353.5

OS 15679g PTMC 3-P-2 Potomac River 80–82 Mulinia 0.01 540 30 1860–1900 AD 86.5F10.8 453.5

WW 1291g PTXT 2-G-2 Patuxent River 85–87 Mulinia 0 510 60 1880–1910 AD 86.5F10.8 423.5

WW 1292g PTXT 2-G-2 Patuxent River 95–97 Mulinia 0 520 50 1880–1910 AD 86.5F10.8 433.5

WW 2703g LCPTK 1-P-3 Little Choptank River 360–362 Mulinia 0 460 55 1880–1910 AD 86.5F10.8 373.5

WW 2704g LCPTK 1-P-3 Little Choptank River 406–408 Mulinia 0 540 55 1880–1910 AD 86.5F10.8 453.5

OS 19212h RD 98-1 Rhode River 142 shell �0.04 325 60 1880–1910 AD 86.5F10.8 238.5

OS 19216h RD 98-1 Rhode River 203 shell �0.4 325 60 1880–1910 AD 86.5F10.8 238.5

a d13C notation relative to Pee Dee Belemnite standard. Values of 0 are assumed; all others were measured.b One-sigma counting error.c Samples selected at depths identified as representing peak land clearance (1880–1910 AD) based on Ambrosia pollen abundance (see Willard et al., 2003).d Expected 14C age calculated as average of single-year age determinations of wood from 1880 to 1910 AD (Stuiver et al., 1998).e Reservoir effect calculated as difference between 14C age and average expected 14C age.f Dates previously published in Willard and Bernhardt (2004).g Dates previously published in Cronin et al., (2000).h Dates previously published in Colman et al., (2002).

D.A.Willa

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ryChange47(2005)17–35

25


either unconformities or periods of very slow depo-

sition at these sites, with much poorer temporal

resolution than in the remainder of the record.

Because no individual core represents the entire

Holocene, we combined records from three cores

(MD99-2207, MD99-2208, and MD99-2209) to pro-

duce a composite series representing the last 10.5 ka.

To avoid the possibility of introducing an incorrect

chronological order of samples by simply merging

data from all four cores, we spliced together the

highest resolution continuous core intervals for each

time period. This resulted in omission of data from

core PTMC 3 in the composite record because of

temporal overlap with core MD99-2209. We also

omitted samples representing the last 300 years, when

Colonial land clearance and subsequent land-use

practices became the primary control on plant com-

munity composition in the eastern United States. In the

composite record, the average sample resolution (Dt)

is 27.7 years and varies from 3 to 210 years. The most

highly resolved portions of the record are from 303 to

2097 years (average Dt=37.5 years), 4139–4880 years

(average Dt=12.8 years), and 5909–10,500 years

(average Dt=19.1 years). This resolution permits

detailed comparison of the terrestrial record of climate

variability with marine records from the North Atlantic

Ocean and provides an exceptionally well-resolved

Holocene pollen record from eastern North America.

3.2. Methodology

Pollen was isolated from estuarine sediments using

standard palynological techniques (Traverse, 1988;

Willard et al., 2003). Initial sample spacing was at 10

cm, with subsequent sampling at 2 cm increments to

maximize temporal resolution. For each sample, one

tablet of Lycopodium spores was added to 5–7 g of

dried sediment for calculation of pollen concentration

(pollen/gram dry sediment). Samples were processed

with HCl and HF to remove carbonates and silicates.

Late Holocene samples (b5 ka) were acetolyzed (1

part sulphuric acid/9 parts acetic anhydride) in a

boiling water bath for 10 min, neutralized, and treated

with 10% KOH for 10 min in a water bath at 70 8C.Early Holocene samples (N5 ka) were treated with

cold 35% nitric acid for 5 min before being heated in a

boiling water bath with 10% KOH for 15 min. After

neutralization, residues were sieved with 150 Am and

10 Am mesh to remove the coarse and clay fractions,

respectively. When necessary, samples were swirled in

a watch glass to remove mineral matter. After staining

with Bismarck Brown, palynomorph residues were

mounted on microscope slides in glycerin jelly. At

least 300 pollen grains were counted from each

sample to determine percent abundance and pollen

concentration. Confidence limits for Pinus and

Quercus percentages were calculated using binomial

standard errors as outlined in Buzas (1990). Mann–

Whitney tests were used to determine whether

abundance of indicator taxa varied significantly both

within and among cores. Pollen data from surface

samples and sediment cores are available from the

North American Pollen Database (NAPD) at the

World Data Center for Paleoclimatology in Boulder,

CO (http://www.ngdc.noaa.gov/paleo/pollen.html)

and at the U.S. Geological Survey Atlantic Estuaries

Project website (http://geology.er.usgs.gov/eespteam/

Atlantic/index.htm).

3.3. Pollen records from Chesapeake Bay

Pinus and Quercus pollen dominate Holocene

records from Chesapeake Bay sediment cores, collec-

tively averaging 79% of pollen assemblages. Carya

and other hardwood taxa contribute minor amounts of

pollen to sediments, as do herbaceous taxa such as

grasses and sedges (Fig. 4). Quercus is a prominent

component in both oak–chestnut and oak–pine–hick-

ory forests, and Mann–Whitney tests show that neither

its percent abundance nor concentration (pollen/gram

dry sediment) varied significantly before Colonial land

clearance in the 17th–18th centuries. After land

clearance, however, oak pollen abundance was nearly

halved (Fig. 4). Pinus, with limited abundance and

distribution in the Chesapeake Bay watershed, is

characterized by significant ( pb0.001, using Mann–

Whitney tests) fluctuations in both percentage and

concentration data between high-pine and low-pine

intervals of the Holocene; decreases during low-pine

intervals are greater than confidence intervals calcu-

lated for the counts (~4.5%: see Fig. 5) (Buzas, 1990).

A striking feature of Chesapeake Bay pollen

records is the doubling of Pinus pollen between

~5.5 and 4.8 ka (Figs. 5 and 6A). The change

represents the mid-Holocene shift from deciduous to

mixed deciduous–conifer forests in the region. The


http://geology.er.usgs.gov/eespteam/atlantic/index.htm

A. MD 99-2207

100200300400500600700800900

10001100120013001400150016001700180019002000

⊗ 490

⊗ 4140

⊗ 4630

⊗ 4810⊗ 5400⊗ 6560⊗ 6900⊗ 7025⊗ 7560

⊗ 9210⊗ 9400

⊗ 10500

0 10 20 30 40 50 60 70 0 100 10 20 30 40 50 60 70 0 0 00

B. MD99-2209

100

200

300

400

500

600

700

800

900

1000

1100

1200

1300

1400

1500

1600

1700

Pinus

Carya

Querc

us

Tsuga

Poace

ae

Cyper

acea

e

Ambr

osia

⊗

⊗

⊗⊗

⊗⊗⊗ 270

3404106757701165

1610181021404230

5905

6100

63806390

6700

673072007570

⊗

⊗

⊗

⊗

⊗⊗

⊗⊗

⊗

⊗

⊗

Cal. yr

BP

Dep

th (

cm)

Dep

th (

cm)

Percent Abundance

Post-ColonialLand

Clearance

LateHolocene

(2.1-0.25 ka)

Condensed Section

EarlyHolocene

(7.2-5.9 ka)

Post-ColonialLand

Clearance

LateHolocene

(5.0-4.1 ka)

Condensed Section

EarlyHolocene

(10.5-5.0 ka)

0

0

b

c

d

a

c

d

Fig. 4. Percent abundance of pollen from major plant taxa in (A) core MD99-2207 and (B) MD99-2209. Both core locations are shown in Fig. 1.

Circles with bxQ indicate radiocarbon dates; filled circles represent age estimates from biostratigraphic horizons (a–d). (a) Quercus increase

~10.5 ka; note that this is a conservative estimate and that basal sediments may be younger. (b) Carya increase ~9.4 ka. (c) Tsuga decline at 5.4

ka. (d) Ambrosia rise ~1880–1910 AD (see text for discussion of biostratigraphic events).


30 40 50 6010 20 30 40 5010 20 30 40 50

PTMC-3M D 9 9 - 2 2 0 9M D 9 9 - 2 2 0 7

10 20 30 40 50

0

100

200

300

400

500

600

700

800

900

1000

1100

1200

1300

1400

1500

1600

1700

1800

1900

2000

2100

M D 9 9 - 2 2 0 8

50067575010602080410

680770

1165

1610

181021404340

5905

6100

63806390

6700

7200

4140

4630

48105400656069007025

7560

9400

<10,500

282

5447

6462

741078108217

2?1

9210

% Pinus

0

Dep

th in

cor

e (c

m)

{{

{

{

{

{

{

{

{4A

3

4A

5

6

0

1

4A

0

*

*

*

*

{270

{

4 {

{

Fig. 5. Percent abundance of pine pollen vs. depth in four Chesapeake Bay cores; all core sites are indicated in Fig. 1. Vertical lines indicate early and late Holocene mean pine

abundance at each site. Numbers 0–6 refer to North Atlantic millennial-scale cold events (Bond et al., 2001). Event 4A represents an additional cool event ~7 ka that was not

recognized in the North Atlantic record. Black dots indicate radiocarbon dates; clear circles indicate biostratigraphic horizons. Lines with asterisks indicate the 95% confidence

intervals (F4%) for the counts.

D.A.Willa

rdet


ryChange47(2005)17–35

28


well-documented migration of southern pines along

the Atlantic Coastal Plain from Florida northward to

Pennsylvania and New Jersey has been attributed to

warmer and wetter late Holocene winters resulting

from orbitally driven solar insolation changes (Watts,

1979; Webb et al, 1987). Using temperature estimates

derived from surface sample calibrations of pollen, the

mid-Holocene shift corresponds to a mid-Atlantic

January warming of up to 2–4 8C from the early to

late Holocene. This is consistent with model recon-

structions of warmer late Holocene winters due to

increased winter insolation and slightly greater pre-

cipitation in eastern North America since 6 ka

(Harrison et al., 2003; Kutzbach et al., 1998).

Superposed upon the mid-Holocene shift are a

series of oscillations in Pinus abundance that suggest

millennial- to centennial-scale temperature variability

(Figs. 5 and 7). Multi-taper method (MTM) spectral

analysis (Thomson, 1982) has been employed to

quantitatively test for centennial- to millennial-scale

periodic components within the composite Pinus

abundance time series. This analysis has been

restricted to the longest uninterrupted interval of

high-resolution pollen data, between 5909 and

10,500 years BP, to avoid the relatively large temporal

gaps (up to 210 years) in portions of the composite

record (Fig. 6A). Between 5909 and 10,500 years BP,

the time series is characterized by sample intervals that

range from 6 to 65 years (average Dt=19.1 years,

standard deviation=9.4 years); these intervals permit

robust quantification of periodicities as short as 130

years (2*65 years). Following linear interpolation to an

even sampling interval (5 years) and removal of a

linear trend, MTM power and harmonic spectra were

calculated using five 3k discrete prolate spheroidal

sequence data tapers (Thomson, 1982).

MTM spectral analysis permits assessment of the

variance contribution (power) at discrete frequencies

in the Pinus abundance time series, and also provides

a statistical F-test for the presence of pure sinusoidal

(harmonic) components in the time series. The results

of the MTM harmonic and power spectral analyses are

displayed in Fig. 6. The harmonic analysis results

indicate five highly significant (N90% significance)

periodic components in the centennial–millennial

band: 1429, 521, 282, 177, and 148 years (Fig. 6C,

gray spectrum). The 1429-year periodic component

displays the greatest amplitude (Fig. 6C, bold line),

and the power spectrum (Fig. 6B, bold line) indicates

that variability within the millennial band is the

dominant source of variance in this portion of the

Pinus abundance record. Relatively high power is also

associated with the significant periodic components

identified at 282 and 148 years.

Millennial-scale periodic variability in the entire

Holocene time series is expressed as relatively

prolonged minima in Pinus abundance (events num-

bered 0–6 in Fig. 7). These minima are clearly

observed following smoothing with a 400-year moving

average (Fig. 7A), and in lowpass-filtered versions of

the Pinus abundance record (Fig. 7B). The spacing of

the centers of the minima is not precisely 1429 years,

but varies due to interference between the identified

periodic components, inaccuracies in the time model,

additional non-periodic noise in the climate proxy

record, and/or temporal variability in the millennial-

scale periodic forcing. Because southern pine distri-

bution is limited primarily by winter temperature, we

interpret these sustained Pinus minima as cooler, drier

intervals in which January temperature decreased by

0.28 to 2 8C, resulting in minor decreases in Pinus

abundance in adjacent forests.

Several of the Holocene pine minima are note-

worthy. Events 6 and 5 (centered at 9.5 and 8.1 ka,

respectively) are preserved in freshwater sediments

deposited in the channel of the paleo-Susquehanna

River before formation of the Chesapeake Bay estuary

by sea-level rise ~7.6 ka. Rapid mean sedimentation

rates in this interval (~0.36 cm year�1) resulted in

high temporal resolution. The termination of Event 5

coincides with a shift from fresh to brackish con-

ditions associated with the final stages of sea-level rise

associated with Laurentide and Antarctic ice sheet

decay (Berke and Cronin, 2004; Vogt et al., 2000). In

the pollen record the corresponding decrease in

riverbank/marsh herb abundance (Cyperaceae, Poa-

ceae) (Fig. 4) represents the impact of transgression

across the paleo-Susquehanna shoreline. Flooding of

the original river channel to form the much wider bay

submerged the original shoreline marshes. The greater

distance from the subsequent shoreline to the core site

apparently was too great for effective transport of

marsh plant pollen to the site, minimizing their

representation in pollen assemblages.

Event 5, lasting from ~8.3 ka to ~8.0 ka,

corresponds to a widespread cool event centered at

10500

9500

8500

7500

6500

5500

4500

3500

2500

1500

500

0 10 20 30 40 50 60

Tim

e (

years

BP

)

% Pinus

0

0.005

0.01

0.015

0.02

0.025

0.03

0.035

0.04

0

1

2

3

4

5

6

0 0.001 0.002 0.003 0.004 0.005 0.006 0.007 0.008

Am

plit

ude

F-t

est (g

ray)

Frequency (cycles/year)

0

0.00005

0.0001

0.00015

Po

we

r

130 yr

Nyquist

Period

(Max.)

1429 yr 521 yr 282 yr 148 yr177 yr

95% Significance

90% Significance

Resol.

CompositePollen Record

A)

B)

C)

SP

EC

TR

AL A

NA

LY

SIS

CO

RE

2208

2209

2207

2207

2209

Fig. 6. (A) Composite record of % Pinus pollen from three CB cores (MD99-2207, MD99-2208, MD99-2209). The spacing between solid dots

reflects the variable sampling resolution of the composite record. (B) MTM power spectrum for the composite record from 5909 to 10,500 years

BP, employing five 3 14discrete prolate spheroidal sequence data tapers. (C) MTM harmonic analysis results for the composite record from 5909

to 10,500 years BP, employing five 3 14discrete prolate spheroidal sequence data tapers. Amplitude is plotted as a bold line, and the F-test results

are indicated in gray. 90% and 95% significance levels for the F-test results are indicated as dotted lines.


8.2 ka, in which high-latitude temperatures dropped

4–8 8C and marine and terrestrial sites cooled by 1.5–

3 8C (Alley et al., 1997; Barber et al., 1999). Using

the pine–temperature correlation in Fig. 2, Chesa-

peake Bay pollen data suggest a decrease in atmos-

pheric January temperature of up to 3 8C in the mid-

Fig. 7. Composite record of pine pollen abundance from three CB cores

(308N) (Berger and Loutre, 1991) and 14C record from GISP2 ice core (Bo

and a 400-year moving average is superimposed in red. The insolation cu

results of a lowpass filtering of the composite % Pinus record. Data filteri

years (thick line) and 500 years (thin line). A linear trend was removed from

ice-core record of 14C is shown in black. Gray boxes indicate the duratio

Atlantic ice-rafting events (Bond et al., 1999, 2001). The timing of Greenla

(O’Brien et al., 1995), glacial advances (Denton and Karlen, 1973), and co

2000) are shown by bars.

Atlantic United States. Proxy evidence and model

results indicate that the primary driver of this event

was a rapid meltwater release associated with collapse

of the Laurentide ice sheet (Barber et al., 1999; von

Grafenstein et al., 1998; Renssen et al., 2001), but 14C

and 10Be records also indicate a concomitant solar

(MD99-2207, -2208, -2209) compared with Dec. insolation record

nd et al., 2001). In (A), pine abundance is illustrated by a black line,

rve is shown by the dashed black line. In (B), the red line illustrates

ng employed a 20% cosine window, with cutoff frequencies of 1000

the composite % Pinus record prior to lowpass filtering. The GISP2

n of cool mid-Atlantic climate; associated numbers indicate North

nd cold intervals (Johnsen et al., 2001), times of polar cell expansion

oler sub-tropical Atlantic sea-surface temperatures (deMenocal et al.,

0 1 2 3 4 6

205

210

215

220

225

230

10

20

30

40

50

60

Pin

us %

Dec

. Ins

olat

ion

(W m

-2)

-0.2

-0.1

0

0.1

0.2

14C

Res

idua

ls

Glacial advances

NGRIP cold intervals

Polar cell expansion

5

Calendar age in kyrs

0 1 2 3 4 5 6 7 8 9 10

Subtropical Atlantic cool intervals

Pin

us %

(40

0 yr

mov

ing

aver

age)

Tim

ing

of C

ool E

vent

s

A)

B)4A

0 1 2 3 4 5 6 7 8 9 10

Pin

us %

(fil

tere

d, d

etre

nded

)

-0.15

-0.1

0

0.05

0.05

0.1

0.15

Core 2208 22092209 22072207



minimum (Bond et al., 2001). The large decrease in

abundance of pine pollen during this time probably

represents the paired impacts of cooler temperatures

and drier, windier conditions documented from high-

latitude sites (Alley et al., 1997) and the mid-continent

United States (Hu et al., 1999).

Another pine minimum centered at 1.8 ka (Event

1) corresponds to relatively cooler spring conditions

in Chesapeake Bay, documented by Mg/Ca ratios

from foraminiferal shells (Cronin et al., 2003). The

most recent pine minimum associated with the Little

Ice Age (LIA: Event 0) cooling represents a two-step

event (Fig. 5), with the first, more severe, minimum

between 650 years BP and 550 years BP and the

second between 450 years BP and 350 years BP.

During the LIA cooling events, Chesapeake Bay

waters cooled by 2–4 8C (Cronin et al., 2003), and

ice-rafted debris were more abundant in the North

Atlantic Ocean (Bond et al., 1999). Because land

clearance began early in the 17th century and was

well-established by 1750 AD, the subsequent warm-

ing after the LIA is obscured in pollen records by

anthropogenic changes to the regional vegetation.

4. Correlation of terrestrial and marine records

The approximate 1400-year periodicity of mid-

Atlantic pine minima/cool intervals is similar to

periodicities of cold intervals in the North Atlantic

identified using petrographic indicators (Bond et al.,

2001), d18O records of cooler conditions and/or

changes in atmospheric circulation interpreted from

the NGRIP ice core and from Irish speleothems

(Johnsen et al., 2001; McDermott et al., 2001),

foraminiferal evidence for cooling in the sub-tropical

Atlantic Ocean (deMenocal et al., 2000), sea-salt (K)

records of increased aridity from the GISP2 ice core

(O’Brien et al., 1995), and biological and geochem-

ical records from southwestern Alaska (Hu et al.,

2003), even with the inherent limitations of chrono-

logical comparisons among different records (Fig. 7).

The correlations are poorest between ~2.2 ka and 4 ka

and 5 ka and 5.8 ka, when sedimentation rates were

slow or where unconformities exist in these Ches-

apeake Bay cores. In the early and late Holocene,

however, the peaks and troughs match extremely well,

and the quasi-periodicity of ~1400 years is compara-

ble to the pattern shown by North Atlantic drift-ice

records (Bond et al., 1997, 1999, 2001).

The North Atlantic Holocene cold cycles have

been linked in unknown ways to solar variability

through comparison with cosmogenic isotope records

(Bond et al., 2001). A potential explanation for

millennial-scale patterns observed in Holocene proxy

records lies in the relationship between solar wind

strength and large-scale, atmospheric pressure systems

on earth over decadal to centennial time scales

(Angell, 1998; Boberg and Lundstedt, 2002; Tinsley,

2000; Tinsley and Heelis, 1993). Solar wind strength

has been linked to cloudiness, configuration of the

circumpolar vortex and planetary waves, and to

latitudinal shifts of storm tracks across the North

Atlantic Ocean documented between sunspot maxima

and minima (Tinsley, 2000). Solar maxima have been

correlated with increased Northern Hemisphere land

temperature (1861–1989 AD) and decreased sea-ice

extent around Iceland (1740–1970 AD) (Friis-Chris-

tensen and Lassen, 1991). Likewise, using general

circulation models, Shindell et al. (2001) showed that

reduced solar irradiance during the late 17th century

Maunder Minimum altered atmospheric circulation

patterns in the North Atlantic region, resulting in

colder winters over northern hemisphere continents.

Such results are consistent with the hypothesis that

solar variability directly influences atmospheric cir-

culation patterns and climate on a global scale over a

variety of time scales.

The similar periodicities exhibited by Chesapeake

Bay pine minima and cosmogenic isotope (10Be and14C) maxima (Fig. 7B), indicating low solar activity,

suggest long-term, possibly solar-driven shifts toward

a periodically expanded circumpolar vortex and

altered jet stream. Intervals of increased Holocene

Arctic storminess and aridity (Mayewski et al., 1997)

and altered eastern European lake levels (Magny et

al., 2002) have similar periodicities to the Chesapeake

Bay record and have been attributed to fluctuating

polar cell size and resultant changes in the jet stream.

The hypothesis of periodic, long-term solar minima

catalyzing extended intervals of circumpolar vortex

expansion and enhanced meridionality could explain

why no synchronous, ubiquitous cooling responses

have been documented in response to decreased solar

activity during observed millennial-scale year cycles

(Rind, 2002). Changes in the configuration of the


planetary wave pattern are manifested by regional

differences in climate rather than a globally uniform

response (Davis and Benkovic, 1992). Other forcing

factors also may amplify or dampen the regional

response to solar variability. For example, although

long-term records of vulcanism from ice cores do not

share the observed ~1400 periodicity, extensive

vulcanism between 7 and 9 ka and during the Little

Ice Age (Zielinski et al., 1994) may have augmented

climatic extremes related to solar minima. Changes in

oceanic thermohaline circulation (Bond et al., 2001)

also may be related to changes in atmospheric

circulation. As additional terrestrial and marine

records exhibiting similar periodicities are developed,

it should be possible to better understand the relative

influence of solar variability on global, long-term

atmospheric circulation patterns and compare it to

other external factors influencing these millennial-

scale events.

Acknowledgements

We thank the crew and scientific staff of the R/V

Marion Dufresne and the IMAGES V program for

providing cores MD99-2207, -2208, and -2209 and for

other assistance with the cruise. We also thank Rick

Younger and the crew of the R/V Kerhin for assistance

in piston coring. Cheryl Vann provided assistance with

statistical analyses. We acknowledge helpful discus-

sions with Gerard Bond and Thomas Cronin during the

course of the research. Chris Swezey, Blaine Cecil, and

Dorothy Peteet provided helpful comments on earlier

versions of the manuscript. We thank Stephen Pekar

and an anonymous reviewer for subsequent reviews of

the manuscript. We acknowledge the laboratory and

field assistance of P. Buchanan, J. Damon, A.

Fagenholtz, J. Murray, T. Sheehan, N. Waibel, and L.

Weimer. This research was supported by the U.S.

Geological Survey Earth Surface Dynamics and Prior-

ity Ecosystem Studies Programs.

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